The primary atmospheric source of soot or black carbon particles is combustion
of fossil fuels and biomass burning at the Earth's surface, with total emission
values near 12 Tg C yr-1 (Liousse et al., 1996). This value exceeds reasonable
estimates of the aircraft source of black carbon by several orders of magnitude
(Bekki, 1997). For example, aircraft are estimated to have emitted 0.0015 to
0.015 Tg C as soot into the atmosphere in 1992 [with EI(soot) of 0.01 to 0.1
g C/kg fuel] (Friedl, 1997; Rahmes et al., 1998). As in the case of sulfate
aerosol, deposition and scavenging of black carbon near surface sources creates
large vertical gradients in the lower atmosphere, with soot concentrations falling
by 1 to 2 orders of magnitude between the surface and the lower stratosphere
(Penner et al., 1992; Cooke and Wilson, 1996; Liousse et al., 1996). A possible
meteoritic source of soot in the lower stratosphere has been considered but
is not well quantified at present (Chuan and Woods, 1984).

Few direct measurements of soot abundance are available in the upper troposphere
and lower stratosphere. The most extensive measurements in these regions are
from aircraft impactor measurements (Pueschel et al., 1992, 1997; Blake and
Kato, 1995). The accuracy of such measurements depends on knowledge of the impactor
for small soot particles. The results (Figure 3-10)
are considered to represent a lower limit for soot number and mass because of
size-selective sampling and because of scavenging of soot by background aerosol
particles. Features of the measurements include a large gradient between the
Northern and Southern Hemispheres and large variability with altitude at northern
mid-latitudes. The large vertical variability of soot at northern mid-latitudes
cannot be explained by the accumulation of aircraft soot emissions. With typical
soot mass densities observed to be approximately 1 ng m-3, soot is estimated
to represent approximately 0.01% of the stratospheric aerosol mass (Pueschel
et al., 1992). In other sampling flights over southern Germany, measurements
of absorbing mass (probably soot) at 8-12 km altitude and partly within cirrus
clouds showed concentrations above 10 ng m-3, with higher values correlated
with local aviation fuel consumption (Ström and Ohlsson, 1998).

Modeling studies of soot distribution in the upper troposphere and lower stratosphere
differ on the importance of aircraft sources of soot. Some model simulations
suggest that ground-level sources of soot (12 Tg C yr-1) are as important
as aircraft sources in the upper troposphere and predict maximum soot values
there [2 to 5 ng m-3 in Liousse et al. (1996) and 10 to 50 ng m-3
in Cooke and Wilson (1996)] that exceed observed values near 200 hPa at northern
mid-latitudes (Pueschel et al., 1997). Other model results show that current
aircraft could be a noticeable source of the soot near the tropopause at northern
mid-latitudes (Bekki, 1997; Danilin et al., 1998; Rahmes et al., 1998). Tracer
simulation results from the AER 2-D model (Section 3.5.1)
were multiplied by EI(soot) of 0.04 g/kg fuel (Döpelheuer, 1997) to estimate
the global distribution of soot. The results (Figure
3-11) show maximum values of 0.6 ng m-3 near 12 km at northern
mid-latitudes. These values are in the middle of the range of other model results
and are in the range of observed values (Figure 3-10).
However, because the fleet-mean EI(soot) is uncertain and may range from 0.01
to 0.1 g/kg fuel, the effective range of the maximum in Figure
3-11 is 0.15 to 1.5 ng m-3. At 20 km, fuel tracer simulations
show aircraft-induced zonal mean soot perturbations to be approximately 100
times smaller than maximum observed values (Figure
3-10).

Observations of soot in the upper troposphere and lower stratosphere are too
limited to provide an estimate of any long-term changes in soot concentrations
in those regions. The possible consequences of heterogeneous reactions on soot
(Bekki, 1997; Lary et al., 1997) are discussed in Chapter
2.

3.3.6. Polar Stratospheric Clouds and Aircraft Emissions

During winter in the polar regions, low temperatures lead to the formation
of polar stratospheric cloud (PSC) particles, which contain H2SO4, HNO3, and
H2O (e.g., WMO, 1995; Carslaw et al., 1997; Peter, 1997). PSCs activate chlorine,
leading to significant seasonal ozone losses in the lower stratosphere, particularly
in the Southern Hemisphere (WMO, 1995). PSC formation may be enhanced by the
atmospheric accumulation of aircraft emissions of NOx, H2O, and sulfate, as
well as through direct formation in aircraft plumes in polar regions (Section
3.2 and Chapter 4). If aircraft emissions change
the frequency, abundance, or composition of PSCs, the associated ozone loss
may also be modified (Peter et al., 1991; Arnold et al., 1992; Considine et
al., 1994; Tie et al., 1996; Del Negro et al., 1997). The effects of subsonic
aircraft emissions on PSCs and stratospheric ozone are expected to be smaller
than those of similar emissions from supersonic aircraft because subsonic emissions
occur in the 10- to 12-km region, whereas supersonic emissions will most likely
occur in the 15- to 20-km region. Ambient temperatures in the 10- to 12-km region
are usually too high (> 200 K) for PSCs to form with available H2O and HNO3,
and ozone and total inorganic chlorine concentrations are much lower than near
20 km.

The impact of the subsonic fleet on PSC formation has not been well studied.
The results of the fuel tracer simulation discussed in Section
3.3.4 can be used to estimate the increase of PSC SAD as a result of aircraft
emissions of H2O and NOx. Assuming an EI(H2O) of 1,230 g/kg, EI(NOx) of 15 g/kg,
complete conversion of NOx to HNO3, and formation of NAT particles at threshold
temperatures, AER model results for a 1992 subsonic fleet show an additional
condensation of HNO3 on NAT particles ranging from 0.02 ppbv at 60°N to 0.12
ppbv at 85°N at 20 km in January. These values provide an increase of 0.08 mm2
cm-3 in PSC SAD at 20 km and 85°N, assuming a unimodal distribution of PSC particles
with diameter of 1 mm. The increase in spatial extent of PSCs both vertically
and latitudinally is small in the model. PSC increases are very sensitive to
background temperature, H2O, and HNO3 values and will differ considerably among
models. The increases are not likely to significantly alter ozone changes in
polar winter because the SAD increases are much less than typical values of
1-10 mm2 cm-3 calculated for PSC events, and satellite data observations show
that the probability of PSC formation below 14 km in the Arctic is generally
very low (< 1%) (Poole and Pitts, 1994).

An important caveat related to the assessment of additional PSC formation as
a result of aircraft emissions is that plume processes are not included. Global
models generally assume that aircraft emissions are homogeneously distributed
in a model grid box that is much larger than an aircraft plume. The consequences
of this assumption have not yet been fully evaluated. In one model study, reactions
on PSCs did not affect ozone chemistry in a subsonic plume at northern mid-latitudes
in April (Danilin et al., 1994). A further caveat is that estimated PSC changes
from aircraft emissions have not accounted for projected cooling of the stratosphere,
which may enhance PSC formation.

The chemical implications of increased PSC formation for ozone chemistry and
atmospheric composition are further discussed in Chapter 2.
The effects of future aircraft fleets on additional PSC formation and subsequent
ozone response are presented in Chapter 4.